2
Emissions, Concentrations, and Related Factors

2.1CONTRIBUTION OF DIFFERENT CHEMICALS TO CO2EQUIVALENT LEVELS AND CLIMATE CHANGES

A range of anthropogenic chemical compounds contribute to changing Earth’s energy budget, thereby causing the planet’s global climate to change. For example, increases in greenhouse gases absorb infrared energy that would otherwise escape to space, acting to warm the planet, while some types of aerosol particles can contribute to cooling the planet by reflecting incoming visible light from the Sun. These components of our atmosphere are emitted from a variety of human activities, including for example fossil fuel burning, land-use change, industrial processes such as cement production, and agriculture. The gases and particles involved are frequently referred to as drivers of climate change, or radiative forcing agents. Detailed reviews of radiative forcing is presented in Forster et al. (2007) and Denman et al. (2007). Radiative forcing due to various climate change agents can be converted to equivalency with the concentration of CO2 (CO2 equivalent), one frame of reference for this report (see Figure 2.1). Here we briefly summarize how major forcing agents contribute to current and future CO2-equivalent target levels and explore implications for global mean temperature increases.

Some greenhouse gases and aerosols are retained for days to years in the atmosphere after emission. The concentrations of such compounds in the atmosphere are tightly coupled to the rate of emission. Their concentrations would drop rapidly if emissions were to cease. Increasing emissions lead to increases in concentrations of such gases, while constant emissions are required for their concentrations to be stabilized. Methane is a key greenhouse gas with an atmospheric lifetime of about 10 years whose concentration has approximately doubled since the pre-industrial era (1750), and it is the second most important greenhouse gas, currently contributing about

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2
Emissions, Concentrations,
and Related Factors
2.1 CONTRIBUTION OF DIFFERENT CHEMICALS TO CO2
EQUIVALENT LEVELS AND CLIMATE CHANGES
A range of anthropogenic chemical compounds contribute to chang-
ing Earth’s energy budget, thereby causing the planet’s global climate to
change. For example, increases in greenhouse gases absorb infrared energy
that would otherwise escape to space, acting to warm the planet, while
some types of aerosol particles can contribute to cooling the planet by
reflecting incoming visible light from the Sun. These components of our
atmosphere are emitted from a variety of human activities, including for
example fossil fuel burning, land-use change, industrial processes such
as cement production, and agriculture. The gases and particles involved
are frequently referred to as drivers of climate change, or radiative forcing
agents. Detailed reviews of radiative forcing is presented in Forster et al.
(2007) and Denman et al. (2007). Radiative forcing due to various climate
change agents can be converted to equivalency with the concentration of
CO2 (CO2 equivalent), one frame of reference for this report (see Figure 2.1).
Here we briefly summarize how major forcing agents contribute to current
and future CO2-equivalent target levels and explore implications for global
mean temperature increases.
Some greenhouse gases and aerosols are retained for days to years in
the atmosphere after emission. The concentrations of such compounds in the
atmosphere are tightly coupled to the rate of emission. Their concentrations
would drop rapidly if emissions were to cease. Increasing emissions lead
to increases in concentrations of such gases, while constant emissions are
required for their concentrations to be stabilized. Methane is a key green-
house gas with an atmospheric lifetime of about 10 years whose concentra-
tion has approximately doubled since the pre-industrial era (1750), and it
is the second most important greenhouse gas, currently contributing about
59

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60 CLIMATE STABILIZATION TARGETS
25 ppmv of CO2 equivalent (see Figure 2.1). From about 1998 to 2007,
methane concentrations remained nearly constant (Forster et al., 2007).
However, methane began to increase after 2007. In the absence of mitiga-
tion, methane is expected to continue to make significant contributions to
climate change during the 21st century (see Section 2.2).
FIGURE 2.1 (left) Best estimates and very likely uncertainty ranges for aerosols and gas contributions to
CO2-equivalent concentrations for 2005, based on the radiative forcing given in Forster et al. (2007). All
major gases contributing more than 0.15 W m–2 are shown. Halocarbons including chlorofluorocarbons,
hydrochlorofluorocarbons, hydrofluorocarbons, and perfluorocarbons have been grouped. Direct effects of
all aerosols have been grouped together with their indirect effects on clouds. (right) Total CO2-equivalent
concentrations in 2005 for CO2 only, for CO2 plus all gases, and for CO2 plus gases plus aerosols.

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EMISSIONS, CONCENTRATIONS, AND RELATED FACTORS 61
In sharp contrast, some greenhouse gases have biogeochemical proper-
ties that lead to atmospheric retention times (lifetimes) of centuries or even
millennia. These gases can accumulate in the atmosphere whenever emis-
sions exceed the slow rate of their loss, and concentrations would remain
elevated (and influence climate) for time scales of many years even in the
complete absence of further emission. Like the water in a bathtub, concen-
trations of carbon dioxide are building up because the anthropogenic source
substantially exceeds the natural net sink. Even if human emissions were
to be kept constant at current levels, concentrations would still increase,
just as the water in a bathtub does when the water comes in faster than it
can flow out the drain. The removal of anthropogenic carbon dioxide from
the atmosphere involves multiple loss mechanisms, spanning the biosphere
and ocean (see Section 2.4), and carbon dioxide removal cannot be char-
acterized by any single lifetime. Although some carbon dioxide would be
lost rapidly to the terrestrial biosphere and to the shallow ocean if human
emissions cease, some of the enhanced anthropogenic carbon will remain
in the atmosphere for more than 1,000 years, influencing global climate
(Archer and Brovkin, 2008). The warming induced by added carbon dioxide
is expected to be nearly irreversible for at least 1,000 years (Matthews and
Caldeira, 2008; Solomon et al., 2009), see Section 3.4.
Figure 2.1 shows that carbon dioxide is the largest driver of current
anthropogenic climate change. Other gases such as methane, nitrous oxide,
and halocarbons also make significant contributions to the current total
CO2-equivalent concentration, while aerosols (see Section 2.3) exert an
important cooling effect that offsets some of the warming. The best estimate
of net total CO2 equivalent concentration of the sum across these forcing
agents in the year 2005 is about 390 ppmv (with a very likely range from
305 to 430 ppmv). Global carbon dioxide emissions have been increasing at
a rate of several percent per year (Raupach et al., 2007). If there were to be
no efforts to mitigate its emission growth rate, scenario studies suggest that
carbon dioxide could top 1,000 ppmv by the end of the 21st century. Carbon
dioxide alone accounts for about 55% of the current total CO2-equivalent
concentration of the sum of all greenhouse gases, and it will increase to
between 75 and 85% by the end of this century based on a range of future
emission scenarios (see Section 2.2). Thus carbon dioxide is the main forcing
agent in all of the stabilization targets discussed here, but the contributions
of other gases and aerosols to the total CO2-equivalent remain significant,
motivating their consideration in analysis of stabilization issues.
How large a reduction of emissions is required to stabilize carbon di-
oxide concentrations, and does it depend upon when it is done or on the

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62 CLIMATE STABILIZATION TARGETS
chosen target stabilization concentration? Studies over the past five years
of so using many different carbon cycle models have improved our under-
standing of requirements for carbon dioxide stabilization. This is because of
more detailed treatments of carbon-climate feedbacks, including the ways
in which warming decreases the efficiency of carbon sinks as compared to
earlier work (e.g., Jones et al., 2006; Matthews, 2006). Figure 2.2 shows
an example of stabilization for two different Earth Models of Intermediate
Complexity (EMICs), the University of Victoria (UVIC) model and the Bern
model (see Methods section for descriptions of these two models; see also
Plattner et al., 2008, and references therein for a model intercomparison
study). In this example test case, carbon dioxide emissions increase at cur-
rent growth rates of about 2% per year to a maximum of about 12 GtC per
FIGURE 2.2 Illustrative calculations showing CO2 concentrations and related warming in two EMICS (the
Bern model and the University of Victoria model, see Methods) for a test case in which emissions first in-
crease, followed by a decrease in emission rate of 3% per year to a value 50%, 80%, or 100% below the peak.
The test case with 100% emission reduction has 1 trillion tonnes of total emission and is also discussed in
Section 3.4.

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EMISSIONS, CONCENTRATIONS, AND RELATED FACTORS 63
year, followed by a decrease of 3% per year down to a selected total reduc-
tion of 50, 80, or 100%. The rate of decrease of 3% per year used here is
derived from scenario analysis described in the next section. This section
together with the next section aim to probe what plausible rates of emis-
sions reduction based upon scenario studies imply for the future evolution
of carbon dioxide concentrations. The rate of possible emissions reductions
of carbon dioxide depends upon factors including e.g., commitments to
existing infrastructure and development of alternatives, see Section 2.2. It is
interesting to note that even in the case of the phaseout of ozone-depleting
substances under the Montreal Protocol, emissions reductions were about
10% per year initially but stalled at a total reduction of about 80% of the
peak, with some continuing emissions of certain gases occurring due for
example, the challenge of finding alternatives for fire-fighting applications
(see IPCC, 2005).
Figure 2.2 shows that carbon emission reductions of 50% do not lead to
long-term stabilization of carbon dioxide, nor of climate, in either of these
models, as has also been shown in previous studies (e.g., Weaver et al.,
2007). It is noteworthy that the Bern model has weaker carbon-climate
feedbacks than the UVIC model; nevertheless both models show the need
for emissions reductions of at least 80% for carbon dioxide stabilization
even for a few decades, while longer-term stabilization requires nearly
100% reduction. Very similar results were obtained in other test cases run
for this study considering peaking at higher values, or decreasing at rates
from 1 to 4% per year (see also Meehl et al., 2007; Weaver et al., 2007).
Figure 2.3 shows sample calculations evaluated in Meehl et al. (2007) us-
ing three different models for various stabilization levels. Figure 2.3 shows
that stabilization levels of 450, 550, 750, or 1,000 ppmv require eventual
emission reductions of 80% or more (relative to whatever peak emission
occurs) in all of the models evaluated. Thus current representations of the
carbon cycle and carbon-climate feedbacks show that anthropogenic emis-
sions must approach zero eventually if carbon dioxide concentrations are
to be stabilized in the long term (Matthews and Caldeira, 2008). This is a
fundamental physical property of the carbon cycle and is independent of the
emission pathway or selected carbon dioxide stabilization target. Box 2.1
discusses how emissions of non-CO2 greenhouse gases could affect attain-
ment of stabilization targets.
Figures 2.2 and 2.3 illustrate a fundamental change in understanding
stabilization of climate change that has been prompted by the scientific
literature of the past two years or so (see Jones et al., 2006; Matthews and
Caldeira, 2008). Early work on stabilization using relatively simple models

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64 CLIMATE STABILIZATION TARGETS
FIGURE 2.3 (a) Illustrative atmospheric CO2 stabilization scenarios for 1,000, 750, 550, and 450 ppmv;
SP1000 (red), SP750 (blue), SP550 (green) and SP450 (black), from Meehl et al. (2007). (b) Compatible an-
nual emissions calculated by three models, the Hadley simple model (solid), the UVIC EMIC (dashed), and
the BERN2.5CC EMIC (triangles) for the three stabilization scenarios. Panel (b) shows emissions required for
stabilization without accounting for the impact of climate on the carbon cycle, while panel (c) included the
climate impact on the carbon cycle, showing that emission reductions in excess of 80% (relative to peak val-
ues) are required for stabilization of carbon dioxide concentrations at any of these target concentrations.
suggested that slow reductions in emissions could lead to eventual stabili-
zation of climate (e.g., Wigley et al., 1996). But recent studies using more
detailed models of key feedbacks in the ocean, biosphere, and cryosphere,
have underscored that although a quasi-equilibrium may be reached for a
limited time in some models for some scenarios, stabilizing radiative forcing
at a given concentration does not lead to a stable climate in the long run.
Cumulative emitted carbon can more readily be linked to climate stabiliza-

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EMISSIONS, CONCENTRATIONS, AND RELATED FACTORS 65
BOX 2.1 STABILIZATION AND NON-CO2 GREENHOUSE GASES
Because carbon emissions reductions of more than 80% are required to stabilize carbon dioxide
concentrations, small continuing emissions of carbon dioxide, or emissions of CO2-equivalent
through other gases, could have surprisingly important implications for stabilizing climate
change. For example, emissions of the hydrofluorocarbons (HFCs) currently used as substitutes
for chlorofluorocarbons make a small contribution to today’s climate change. However, because
emissions of these gases are expected to grow in the future if they are not mitigated, and be-
cause of the stringency of the requirement of near zero emissions of CO2-equivalent, these gases
could represent a significant future impediment to stabilization efforts. For example, Figure 2.4
below shows that in the absence of mitigation, the HFCs could represent as much as one-third
of the allowable CO2-equivalent emissions in 2050 required for a stabilization target of 450 CO2-
equivalent. Thus, the analysis presented here underscores that stabilization of climate change
requires consideration of the full range of greenhouse gases and aerosols, and of the full suite
of emitting sectors, applications, and nations.
FIGURE 2.4 Global CO2 and HFC emissions ex-
pressed as CO2-equivalent emissions per year for
the period 2000-2050. The emissions of individual
HFCs are multiplied by their respective GWPs (di-
rect, 100-year time horizon) to obtain aggregate
emissions across all HFCs expressed as equivalent
GtCO2 per year. High and low estimated ranges
based on analysis of likely demand for these
gases and assuming no mitigation of HFCs are
shown. HFC emissions are compared to emissions
for the range of SRES CO2 scenarios, and two 450-
and 550-ppm CO2 stabilization scenarios. The es-
timated CO2-equivalent emissions due to HFCs in
the absence of mitigation reach about 6 GtCO2-
equivalent in 2050, or about a third of the emis-
sions due to CO2 itself at that time in the 450
ppm stabilization scenario. Source: Velders et al.
(2009).
2-4 large 3.eps
bitmap
tion, due to the irreversible character of the induced warming driven by
carbon dioxide (see Section 3.4).
2.2 INFORMATION FROM SCENARIOS
Figure 2.5 shows the emissions of manmade greenhouse gases from
various sectors of the U.S. economy (U.S. EPA, 2008). For highly industri-

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66 CLIMATE STABILIZATION TARGETS
Residential
5%
Commercial
6%
Agriculture
8%
Electric Power
34%
Industry
19%
Transportation
28%
FIGURE 2.5 U.S. greenhouse gas emissions by sector in 2006. Source: U.S. EPA (2008).
2-5.eps
alized countries such as the United States, the difficulty in reducing emis-
sions will depend in large part on the lifetimes of the existing capital stock
associated with the major emitting sectors. The electric sector is the largest
source of manmade emissions in the United States, primarily due to the
carbon dioxide emitted during the combustion of fossil fuels. The lifetime
of coal-fired power plants is measured in decades. The next largest source
of U. S. greenhouse gases is the transportation sector, again due to the
combustion of fossil fuels. Here the lifetime of the capital stock is typically
a decade or two.
Although developed countries historically have been the major emitter
of greenhouse gases, developing countries are on track to overtake them
in the next few years. In their case, the issue becomes one of the capital
stock put in place in the future to support their industrialization process.
With the huge economic growth projected for developing countries and in
the absence of incentives to act otherwise, these countries will likely turn
to the cheapest energy sources to fuel their growth. These fuels currently
are fossil based: coal, oil, and gas. A recent study by the Energy Modeling
Forum, based on eight Energy-Economy models, projected an annual growth
rate of CO2 emissions globally from the burning of fossil fuels and industrial
uses to be of the order of 1 to 2 percent per year over the remainder of the

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EMISSIONS, CONCENTRATIONS, AND RELATED FACTORS 67
century, in the absence of intervention (EMF 22, 2009). The study attributes
much of the growth to developing countries.
Even if wealthier countries like the United States were to reduce their
emissions to zero immediately, it is unlikely that global CO2 emissions
would be stabilized, much less global atmospheric concentrations (Blanford
et al., 2009). Being in their post-industrial phase of development, the eco-
nomic growth rates in developed countries are expected to be lower than
those of developing countries and their mix of goods and services less
carbon intensive. The cumulative reductions of developed countries, even
with aggressive emissions reduction programs, are expected to be low when
compared to those of developing countries.
One important contribution that developed countries can make to
global emissions reductions is to develop the technological wherewithal
that would not only be necessary for their own emission reductions, but
also would be essential for developing countries to meet their economic
development goals with affordable climate-friendly technologies.
As noted above, both the existing capital stock and that put in place in
the future are critical to understanding the difficulty of transitioning away
from the current path of growth in greenhouse gas emissions. Figure 2.6
shows representative carbon pathways (RCPs) for limiting radiative forcing
(watts per m2) at two alternative levels. These are referred to as the RCP
2.61 and RCP 4.5 scenarios. These are among a suite of pathways being
developed for use in the IPCC 5th Assessment. The pathways shown in the
figure were developed by the IMAGE and MiniCAM models, respectively
(Moss et al., 2010).
Figure 2.6 highlights the importance of the carbon budget, that is, the
area under the allowable emissions curve associated with a particular radia-
tive forcing target. Being much lower in the RCP 2.6 scenario than in the RCP
4.5 scenario, we see the rate of growth first slows and then rapidly decline
beginning in 2020. In the case of the higher CO2 budget, emissions rise for
another two decades before peaking. Notice that the maximum rate of de-
cline is comparable in the two scenarios (about 3.5% per year); however, in
the latter it is shifted out in time. The reasons for this shift are both the higher
1Although Moss et al. (2010) refers to this as the RCP 2.6 scenario, this is the one RCP sce-
nario that peaks and then declines. For this reason it is also referred to as the RC P3-PD sce-
nario. The RCP 3-PD has a unique shape. The radiative forcing of RCP 3-PD peaks and declines
(PD), while the radiative forcing of the other RCPs stabilize or rise towards their higher 2100
levels. Specifically, the final RCP 3-PD prepared for climate modeling peaks at 2.99 W/m2 in
2050 and then declines to 2.71 W/m2 in 2100 with the decline continuing beyond 2100. The
decline is due to the availability later in the century of a negative-emitting technology, biomass
with carbon capture and storage (BECs).

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68 CLIMATE STABILIZATION TARGETS
Billion Tons of Carbon
Decline rate 3.5% per year
FIGURE 2.6 shows representative carbon pathways (RCPs) for limiting radiative forcing (W per m2) at two
2-6.eps
alternative levels. The tighter the limit, the earlier the reductions must take effect. With the RCP 2.6 scenario,
the rate of growth first slows and then rapidly declines beginning in 2020. In the case of the less stringent
bitmap
constraint, emissions rise for another two decades before peaking. Here the decline is shifted out in time.
carbon budget and a greater array of low-carbon, economically competitive
alternatives, which are assumed to become available in the future.
We stress that there is a great deal of flexibility regarding the rate at
which new technologies are substituted for existing ones, both on the sup-
ply and demand sides of the energy sector. The rate of retirement of existing
carbon-intensive plant and equipment and their replacement with more
climate-friendly alternatives will depend upon a number of factors. These
include the stabilization target, reference case emissions in the absence
of a price on CO2 (either explicit or implicit), the availability and costs of
alternatives, and the willingness to pay the costs of the transition to a low-
carbon economy. The latter will depend on society’s perception of the ben-
efits (reduction in damages due to climate change). From a purely physical
perspective, decline rates much higher than those shown here are feasible.
It is a matter of the perceived urgency and the motivation to decarbonize.

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EMISSIONS, CONCENTRATIONS, AND RELATED FACTORS 69
Figure 2.7 shows the CO2-equivalent concentrations for these two sce-
narios at three points in time. Notice that in the case of the tighter radiative
forcing goal, there is some “overshoot”. That is, the target is exceeded in the
middle part of the century and then gradually approached. This is due to the
assumption that there will be a “negative” emitting technology, Bioenergy
with Carbon Capture and Sequestration (BECS). Otherwise a faster decline
rate of the capital stock would be required.
2.3 SHORT-LIVED RADIATIVE FORCING AGENTS:
PEAK TRIMMING VERSUS BUYING TIME
The role of CO2 emissions in Earth’s climate future is unique among the
major radiative forcing agents, because the impact of the CO2 emitted into
the atmosphere will continue to alter Earth’s energy budget for millennia to
come. As noted above, Earth’s energy budget is also subject to the influence
of a number of short-lived radiative forcing agents whose radiative effect
would decay to zero on a time scale of weeks to decades if their sources
where shut off. These agents include aerosols, black carbon on snow or
ice, and methane.
Aerosols are produced by burning biomass and fossil fuels, but unlike
CO2 they are not an inevitable by-product of combustion. Some aerosols
reflect energy to space, but other aerosols such as black carbon absorb sun-
light. Reflecting aerosols unambiguously lead to cooling of the surface. Their
effect has offset a portion of the radiative forcing from the anthropogenic
increase of greenhouse gases so far, and any action that reduces reflecting
aerosol emissions will lead to a nearly immediate warming. The effect of
airborne absorbing aerosols is more subtle, because they primarily act to
shift the absorption of solar radiation from the surface to the interior of the at-
mosphere, leaving the top-of-atmosphere energy budget largely unchanged
(Randles and Ramaswamy, 2008). The global mean effect of surface black
carbon is difficult to quantify but is unambiguously a warming, amplified
further by the albedo feedback of melting snow and ice (Flanner et al., 2007;
Hansen and Nazarenko, 2004; McConnell et al., 2007). Aerosol effects can
include direct damage to human health and agriculture, implying that they
should be cleaned up for reasons independent of climate (Agrawal et al.,
2008; Ramanathan et al., 2008). A key question is whether the effort to do
so will help or hurt other efforts to keep warming in check. Although the
discussion of aerosol effects will be based primarily on temperature changes,
it should be kept in mind that the spatially inhomogeneous radiative forcing
from aerosols can lead to regional effects such as changes in clouds and the

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72 CLIMATE STABILIZATION TARGETS
hydrological cycle. These changes constitute an anthropogenic imprint on
climate that is distinct from that associated with the overall warming due
to greenhouse gases.
Methane is currently emitted by a diverse set of anthropogenic sources,
many of which are related to agriculture. Once in the atmosphere, methane
oxidizes to CO2 with a time scale of about a decade. Since one molecule of
CO2 has much less radiative effect than a molecule of methane, for methane
concentrations in the present low range, the resulting CO2 is negligible in
comparison to the CO2 emitted by deforestation and fossil fuel burning. In
contrast to CO2, the climate effect of current anthropogenic methane emis-
sions would decay relatively rapidly if emissions were to cease, as depicted
in Figure 2.8. But it seems difficult to bring anthropogenic methane emis-
sions to zero in the long term, given the continuing need for agriculture to
feed the world’s population. A continued long-term warming contribution
from methane should therefore be anticipated, not because of persistence of
methane in the atmosphere, but because of likely persistence of the source.
For the much more massive methane release that could come from clathrate
destabilization (see Section 6.1), the CO2 produced by oxidation could have
an important effect on climate.
The climate effects of short-lived radiative forcing agents are thus more
reversible than those of CO2, and therefore actions reducing emissions of
short-lived agents have different implications for Earth’s climate future than
actions that affect CO2 emissions. Insofar as it is perceived that control of
methane or black carbon may be technically easier or less economically dis-
ruptive than controlling CO2 emissions, mitigation of the short-lived warm-
ing influences has sometimes been thought of as a way of “buying time” to
put CO2 emission controls into place. This is a fallacy. While one does buy
a rapid reduction by reducing methane or black carbon emissions, this has
little or no effect on the long-term climate, which is essentially controlled
by CO2 emissions because of the persistence of CO2 in the atmosphere. The
situation is illustrated schematically in Figure 2.8. The time course of warm-
ing produced by CO2 emissions alone is given schematically by the black
line. If one adds short-lived radiative forcing agents with an aggregate warm-
ing effect into the mix, the effect will be to add to the temperature increase
until such time as the emissions are brought under control, where after the
temperature will quickly drop back to the CO2-only curve (the blue and red
solid lines on the curve, representing early or delayed mitigation of short-
lived forcing agents). The effect of mitigation of methane and black carbon
is thus to trim the peak warming rather than limit the long-term warming to
which Earth is subjected. If the early action to mitigate methane emissions

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EMISSIONS, CONCENTRATIONS, AND RELATED FACTORS 73
FIGURE 2.8 Qualitative sketch of the time-course of future temperature under various scenarios for control
of emissions of short-lived radiative forcing agents. The time axis is given as time since the beginning of
significant anthropogenic emissions of greenhouse gases. It is assumed that CO2 emissions are brought
to zero after 200 years. SLPF refers to short-lived positive forcing agents, like methane or black carbon on
snow or ice. SLNF refers to short-lived negative forcing agents, primarily aerosols. “Early Action” refers to
a scenario in which early, aggressive action is taken to mitigate emission of short-lived radiative forcing
agents, while “Deferred Action” refers to a scenario in which such actions are delayed. The green line shows
what happens if the aggregate of all short-lived forcings brought under control originally added up to a
cooling effect (so that reducing them warms the climate). The dashed green line is similar, except that it
assumes there is a residual methane emission that cannot be reduced to zero. The cumulative CO2 emis-
sions are assumed to be the same in all of these scenarios.
was done instead of actions that could have reduced net cumulative car-
bon emissions, the long-term CO2 concentration would be increased as a
consequence. Peak trimming in that case would come at the expense of an
increased warming that will persist for millennia. Carbon emission control
and short-term forcing agent control are two separate control knobs that
affect entirely distinct aspects of Earth’s climate and should not be viewed
as substituting for one another.
It would be unrealistic to contemplate policies that would reduce black
carbon emissions while leaving reflecting aerosol emissions intact, given that
the diverse sources of emission yield an interlinked stew of absorbing and

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74 CLIMATE STABILIZATION TARGETS
reflecting aerosols (Ramanathan et al., 2008). The green curve in Figure 2.8
shows what happens if the aggregate of all aerosols brought under control
sums to a cooling effect before mitigation; the mitigation in this case ac-
celerates the approach to the CO2-only curve, as the masking effect of the
aerosols is eliminated. If the long-term situation instead includes a recalci-
trant methane emission rate that is stabilized but not brought to zero, then
the long-term warming is brought above the CO2-only case for a period as
long as the methane emissions continue.
2.4 CARBON CYCLE
The evolution of atmospheric carbon dioxide concentrations depends
on the balance of human emissions, natural processes that remove excess
carbon dioxide from the atmosphere, and the sensitivity of land and ocean
carbon reservoirs to climate change and land use (see Box 2.2). The histori-
cal atmospheric growth rate of carbon dioxide is well-constrained for the
past 50 years from direct instrumental measurements and for periods prior to
that from measurements of gases trapped in ice cores. Global average atmo-
spheric CO2 has risen from a pre-industrial level of about 280 ppm to about
390 ppm by the beginning of 2010. A definitive anthropogenic origin for the
excess carbon dioxide can be assigned based on contemporaneous changes
in carbon isotopes, a parallel decrease in atmospheric oxygen, and by the
fact that the atmospheric carbon dioxide levels for the preceding several
millennia of the Holocene had hovered within plus or minus 5 ppm of the
pre-industrial value. Past fossil fuel combustion rates and carbon emissions
from cement production and land-use change (e.g., deforestation, shifting
land into pasture and agriculture) can be reconstructed, and net land and
ocean carbon sources and sinks can be quantified from a combination of
observations and numerical models. Figure 2.9 presents a recent synthesis
for the global carbon system showing the fluxes between the atmosphere and
various reservoirs versus time (Le Quéré et al., 2009). The terrestrial carbon
fluxes are partitioned with carbon emissions from direct human land-use
change including recovery from earlier human land use, separated from
terrestrial carbon sinks in response to elevated CO2 and climate.
The response of the global carbon cycle to human perturbations can
be characterized by the airborne CO2 fraction, the fraction of the cumula-
tive carbon dioxide emitted by fossil fuel combustion and land-use change
that remains in the atmosphere. The contemporary airborne fraction is cur-
rently slightly less than half (~0.45), and for any specified carbon emission
trajectory, future atmospheric carbon dioxide concentrations depend on the

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EMISSIONS, CONCENTRATIONS, AND RELATED FACTORS 75
BOX 2.2 TIMES CALES FOR REMOVAL OF CO2 FROM THE ATMOSPHERE
The figures show the fate of a pulse of 2,600 Gt of carbon released instantaneously into the
atmosphere as CO2. In the first hundred years CO2 is absorbed into the upper ocean. The resulting
acidification limits further uptake by the upper ocean waters. During this time period, there is also
typically some uptake by the land biosphere. In the next 900 years, the saturated upper ocean
waters mix with the deep ocean, allowing further uptake. Eventually, the deep ocean acidifies as
well, limiting further uptake. Over the next 10,000 years the ocean becomes buffered by dissolu-
tion of carbonate sediments and by carbonates washed in from land, reducing the acidity and
allowing the ocean to take up additional carbon. Over longer time scales spanning more than
100,000 years, most of the remaining CO2 is removed by reacting with silicate minerals to form
carbonates (e.g., limestone). We have not attempted to state the precise time required for silicate
weathering to cause recovery to pre-industrial values, because of uncertainties in silicate weath-
ering parameterization and uncertainties in the long term response of the glacial-interglacial
cycle. The only long-term sink of CO2 is silicate weathering, which is a very slowly increasing
function of temperature. It would require over 20°C of warming to balance a steady state fossil
fuel emission of only a half Gt of carbon per year, so that even an emission as low as this would
lead to a steady accumulation of CO2 in the atmosphere. This calculation does not allow for any
long-term net release of carbon from land ecosystems or marine sediments, though it is known
that the Earth system is capable of such releases. Any such release would increase the long term
CO2 concentrations and delay the recovery to pre-industrial values. (Data up to 10,000 years
based on carbon cycle simulations of Eby et al. (2009). Silicate weathering time scale estimated
from data given in Berner (2004). See Archer et al. (1997), and Archer (2005) for more details on
the mechanisms of CO2 removal.

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76 CLIMATE STABILIZATION TARGETS
FIGURE 2.9 Components of the global CO2 budget. (a) The atmospheric CO2 growth rate. (b) CO2 emissions
from fossil fuel combustion and cement production and from land-use change. (c) Land CO2 sink (negative
Figure 2-9.eps
values correspond to land uptake). (d) Ocean CO2 sink (negative values correspond to ocean uptake). The
land and ocean sinks (c,d) are shown as an duplicate several models normalized ps) observed mean
bitmap (looks like a average of of box 2-9 made from to the
land and ocean sinks for 1990-2000. The shaded area is the uncertainty associated with each component.
Adapted from Le Quéré et al. (2009).

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EMISSIONS, CONCENTRATIONS, AND RELATED FACTORS 77
airborne fraction and how it evolves with time. At present, the ocean and
the land biosphere contribute about equally to the drawdown of excess
carbon dioxide from the atmosphere. Theoretical arguments and numerical
models suggest that the efficiency of both the land and ocean carbon sinks
may decline in the future under warmer climate conditions, which would
act to amplify climate warming (Fung et al., 2005; Friedlingstein et al.,
2006). The magnitude of the climate-carbon cycle feedback, however, var-
ies substantially across model simulations and is a substantial uncertainty
in future climate projections.
Excess atmospheric carbon dioxide dissolves in surface seawater as
inorganic carbon through well-known physical-chemical reactions. The dis-
tribution and global inventory of anthropogenic carbon dioxide in the ocean
are well characterized based on global ship-based observations collected
during the late 1980s and 1990s (Sabine et al., 2004). Ongoing measure-
ments at time-series sites and along ocean sections constrain uptake over
decadal time periods (IOC, 2009). Oceanic anthropogenic CO2 uptake up
to present has been governed primarily by atmospheric CO2 concentrations
and the rate of ocean circulation that exchanges surface waters equilibrated
with elevated CO2 levels with subsurface waters. In particular, key pathways
include the ventilation of the wind-driven thermocline and deep and inter-
mediate water formation. Ocean models constrained by field data provide
estimates of the oceanic transport and air-sea flux of anthropogenic CO2 as
well as reconstructions of past ocean uptake and projections for the future
(Matsumoto et al., 2004; Gruber et al., 2009; Khatiwala et al., 2009).
Future ocean uptake of anthropogenic carbon is expected to decrease
in efficiency (i.e., absorb a smaller fraction of the emissions); the ocean
CO2 sink is expected to continue to increase, but more slowly than the
emissions. Under elevated CO2, the chemical buffer capacity of seawater
decreases, lowering the amount of inorganic carbon absorbed when surface
waters are equilibrated with the atmosphere. Upper-ocean warming reduces
the solubility of carbon dioxide in seawater. Anthropogenic CO2 uptake
will be further reduced because of increased vertical stratification, reduced
ocean ventilation rates, and reduced deep and intermediate water formation
rates, which are expected due to warming in the tropics and subtropics and
increased freshwater input in temperate and polar regions due to elevated
precipitation and sea-ice melt (Sarmiento et al., 1998). In contrast, an in-
crease in the strength of Southern Ocean winds, associated with a more
positive phase of the Southern Annular Mode, may increase future uptake
of anthropogenic CO2 (Russell et al., 2006).
Ocean biogeochemistry plays an important role in ocean carbon stor-

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78 CLIMATE STABILIZATION TARGETS
age, and biogeochemical responses to changing ocean circulation also need
to be considered when assessing future net carbon uptake. The inorganic
carbon concentration in subsurface ocean waters is generally elevated over
surface concentrations because of the downward transport and subsequent
respiration of organic water originally produced in the surface layer. In
coupled carbon-climate models, biogeochemical feedbacks to a warmer
climate tend to partially offset physical-chemical effects and act to reduce
the overall strength of ocean climate-carbon cycle feedbacks. In the South-
ern Ocean, enhanced outgassing of natural CO2 due to stronger winds and
upwelling may more than compensate for increased anthropogenic CO2
uptake, leading to a net reduction in ocean uptake (Le Quéré et al., 2008;
Lovenduski et al., 2007, 2008). Recent observations of the air-sea difference
in the partial pressure of carbon dioxide, the driving force for air-sea CO2
exchange, indicate a weakening of oceanic uptake in a number of regions,
although there remains some debate whether this signal should be attributed
to climate change, ozone depletion, or primarily decadal climate variability
(Le Quéré et al., 2009; Watson et al., 2009).
It is more difficult to directly constrain, on a global scale, the net
fluxes of carbon into and out of the more heterogeneous terrestrial carbon
reservoirs, and terrestrial uptake is often estimated from a combination of
terrestrial biogeochemical models and satellite remote sensing approaches
that have been assessed using process experiments, local CO2 flux towers,
etc. (Canadell et al., 2007; Raupach et al., 2007; Le Quéré et al., 2009).
Land carbon uptake can be computed in a top-down fashion by difference
from the estimated fluxes to the atmosphere, the ocean sink, and the growth
rate in the atmosphere. Slightly more sophisticated approaches utilize the
spatial and temporal variations in atmospheric CO2 with transport models
to infer land and ocean surface fluxes (Rödenbeck et al., 2003; Peylin et al.,
2005). Atmospheric carbon isotope and oxygen/nitrogen ratios also provide
critical constraints on the partitioning of carbon uptake between the ocean
and land biosphere (Rayner et al., 1999).
The contemporary land carbon budget is governed by a combination
of interacting natural and anthropogenic processes rather than any single
mechanism (Pacala et al., 2001; Schimel et al., 2001). Deforestation and
biomass burning result in net CO2 fluxes to the atmosphere as high-carbon
forests are turned into comparatively low-carbon pastures and croplands
(Houghton, 2003). This process is now occurring mainly in the tropics
and is partially countered by temperate regrowth on abandoned farm and
pasture-land (Shevliakova et al., 2009). The impacts of land-use change
can extend for decades after the initial disturbance, and contemporary land

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EMISSIONS, CONCENTRATIONS, AND RELATED FACTORS 79
carbon fluxes may reflect as much aggregate events in the past as current
conditions, complicating the task of deconvolving underlying mechanisms.
Experimentally, fertilization of photosynthesis by higher atmospheric carbon
dioxide levels increases plant growth, in many cases substantially, and con-
tributes to carbon uptake. When included in global models, CO2 fertilization
leads to a modest current sink that may continue for many decades into the
future. Some studies suggest, however, that this effect may be smaller than
earlier thought perhaps due to other limitations such as nitrogen (Long et al.,
2006; Thornton et al., 2009a). Many Northern Hemisphere ecosystems are
also nitrogen limited, and deposition of reactive nitrogen mobilized by fossil
fuel combustion may also cause increased carbon uptake (Lamarque et al.,
2005; Thornton et al., 2009a). Land carbon storage varies on interannual
time scales due to climate variability and in particular variations in tempera-
ture, precipitation, and water availability (Sitch et al., 2008). Wildfire also
plays a major role in the global carbon cycle (Randerson et al., 2006), and
increased tropical fire associated with El Niño droughts may contribute to
increases in the growth rate of atmospheric carbon dioxide concentrations
during recent El Niño years (Van der Werf et al., 2006).
Climate warming may cause land ecosystems to lose carbon because
respiration is more temperature sensitive than photosynthesis, but there is
a wide range of estimates for the climate sensitivity of land carbon stocks.
Larger effects could come if future climates lead to additional disturbances,
especially fire, pests, or widespread replacement of forests to grasslands,
which could rapidly release large amounts of carbon. Current attention
is focusing on both the role of human and climate-caused disturbance in
controlling future ecosystem carbon storage and on physiological processes,
such as carbon dioxide fertilization. Water availability to support photo-
synthesis and primary productivity is thought to be significant, with models
(Fung et al., 2005) and observations (Angert et al., 2005) indicating that
decreasing water balance (drier soil conditions) decreases carbon uptake.
Global effects are a balance between warmer conditions favoring a longer
growing season in the Northern Hemisphere temperate zone, increasing
carbon uptake, and drier soils in the tropics, decreasing carbon uptake (Fung
et al., 2005; Friedlingstein et al., 2006). Past and future land-use change will
also affect land carbon storage and needs to be considered when projecting
future atmospheric CO2 levels.
The combined effects of ocean and land feedbacks have been explored
in a series of coupled climate-carbon models reported by the international
C4MIP (Coupled Climate-Carbon Cycle Model Intercomparison Project;
Friedlingstein et al., 2006; Figure 2.10). Simulations are conducted for the

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80 CLIMATE STABILIZATION TARGETS
20th and 21st centuries forced with specified trajectories for fossil fuel and
land-use CO2 emissions; carbon-climate sensitivity is assessed by comparing
simulations with and without coupling between model atmospheric CO2
and atmospheric radiation (and thus climate). Some simulations exhibit a
large amplifying effect whereby climate effects lead to increased release of
carbon dioxide, mostly from terrestrial processes. By the year 2100, simula-
tions with strong feedbacks have higher atmospheric CO2 levels as much as
200 ppm above the companion simulation without climate feedbacks. The
airborne CO2 fraction tends to increases in models where uptake processes
are more sensitive to climate, where uptake processes are less sensitive to
atmospheric CO2, and where physical climate sensitivity to CO2 is large.
The land modules in the C4MIP simulations did not include interactions
between carbon and other limiting nutrients such as nitrogen. When soil
organic matter is respired due to warming, nitrogen is released, stimulating
enhanced plant growth that can offset the CO2 released from the soil matter.
More recent coupled simulations that include this effect indicate a small
negative climate-carbon feedback on atmospheric CO2 (Thornton et al.,
2009a). Interestingly, the carbon-nitrogen model had only a very weak CO2
fertilization effect on land photosynthesis because of nitrogen limitation.
As a result, the model land carbon uptake responded only weakly to rising
atmospheric CO2 and had as a result one of the largest atmospheric CO2
levels at the end of the 21st century. The degree to which CO2 fertiliza-
tion is modulated by nitrogen is an important unresolved science question,
and for the carbon cycle, sensitivity of the sinks to CO2 is as important as
sensitivity to climate. The C4MIP simulations did not address the full suite
of interactions between the carbon cycle and other changes in the Earth
System; for example, ocean carbon storage can be influenced by ozone-
driven changes in Southern Ocean winds (Le Quéré et al., 2008; Lovenduski
et al., 2008; Lenton et al., 2009). Finally, the current generation of coupled
climate-carbon cycle models neglect a number of carbon reservoirs, such
as high-latitude peats, permafrost and methane clathrates that could be
important on longer time scales (see Section 6.1).